1 2 3 4 5 Smectite formation in the presence of sulfuric acid: Implications for
6 acidic smectite formation on early Mars
7 8 9 10 11 T.S. Peretyazhko1*, P.B. Niles2, B. Sutter1, R. V. Morris2, D. G. Agresti3, L. Le1, D.W. Ming2 12 13 14 15 16 17 18 1 Jacobs, NASA Johnson Space Center, Houston, TX 77058 19 20 2 NASA Johnson Space Center, Houston, TX 77058 21 3 University of Alabama at Birmingham, Birmingham, AL 35294 22 23 24 25 26 27 Corresponding author: T. S. Peretyazhko, Jacobs, NASA Johnson Space Center, 28 Houston, TX 77058 ([email protected]) 29 30 31 32 33 In preparation for
34 Geochimica et Cosmochimica Acta 35 Abstract
36 The excess of orbital detection of smectite deposits compared to carbonate deposits on the martian
37 surface presents an enigma because smectite and carbonate formations are both favored alteration
38 products of basalt under neutral to alkaline conditions. We propose that Mars experienced acidic
39 events caused by sulfuric acid (H2SO4) that permitted phyllosilicate, but inhibited carbonate,
40 formation. To experimentally verify this hypothesis, we report the first synthesis of smectite from
41 Mars-analogue glass-rich basalt simulant (66 wt% glass, 32 wt% olivine, 2 wt% chromite) in the
42 presence of H2SO4 under hydrothermal conditions (~200 °C). Smectites were analyzed by X-ray
43 diffraction, Mössbauer spectroscopy, visible and near-infrared reflectance spectroscopy and
44 electron microprobe to characterize mineralogy and chemical composition. Solution chemistry was
45 determined by Inductively Coupled Plasma Mass Spectrometry. Basalt simulant suspensions in 11-
46 42 mM H2SO4 were acidic with pH ≤ 2 at the beginning of incubation and varied from acidic (pH
47 1.8) to mildly alkaline (pH 8.4) at the end of incubation. Alteration of glass phase during reaction
48 of the basalt simulant with H2SO4 led to formation of the dioctahedral smectite at final pH ~3 and
49 trioctahedral smectite saponite at final pH ~4 and higher. Anhydrite and hematite formed in the
50 final pH range from 1.8 to 8.4 while natroalunite was detected at pH 1.8. Hematite was
51 precipitated as a result of oxidative dissolution of olivine present in Adirondack basalt simulant.
52 Formation of secondary phases, including smectite, resulted in release of variable amounts of Si,
53 Mg, Na and Ca while solubilization of Al and Fe was low. Comparison of mineralogical and
54 solution chemistry data indicated that the type of smectite (i.e., dioctahedral vs trioctahedral) was
55 likely controlled by Mg leaching from altering basalt and substantial Mg loss created favorable
56 conditions for formation of dioctahedral smectite. We present a model for global-scale smectite
57 formation on Mars via acid-sulfate conditions created by the volcanic outgassing of SO2 in the
58 Noachian and early Hesperian.
2
59 1. INTRODUCTION
60
61 Two global eras have been proposed to explain the observed mineralogy on Mars
62 (Bibring et al., 2006). Abundant phyllosilicates dominated by the smectite group (nontronite,
63 montmorillonite and saponite) were formed in the first era under water-rich neutral to alkaline
64 conditions during the Noachian. Formation of sulfate-bearing phases occurred in the second era
65 under acidic conditions likely caused by sulfuric acid during the Hesperian (Poulet et al., 2005;
66 Bibring et al., 2006; Murchie et al., 2009; Bishop et al., 2013). However, such simplified pH-
67 based division of aqueous history conflicts with some mineralogical observations. Large
68 carbonate deposits together with phyllosilicates would be a characteristic of a Noachian Mars
69 dominated by abundant liquid water, neutral/alkaline pH conditions, and a CO2-rich atmosphere
70 (Fairén et al., 2004), but mineralogical observations have detected only isolated carbonate
71 deposits (Ehlmann et al., 2008; Milliken et al., 2009; Morris et al., 2010; Wray et al., 2016). The
72 absence of widespread carbonate deposits could result from carbonate deposition on early Mars
73 followed by its decomposition in acidic environments during Hesperian epoch (Fairén et al.,
74 2004; Bibring et al., 2006; Chevrier et al., 2007). Alternatively, abundant carbonates may never
75 have formed on early Mars because of short-term stability of liquid water and/or lack of dense
76 CO2 atmosphere during Noachian epoch (Bibring et al., 2006; Chevrier et al., 2007; Niles et al.,
77 2013; Edwards and Ehlmann, 2015; Zolotov and Mironenko, 2016). Acidic events could also
78 explain the apparent lack of abundant carbonate/phyllosilicate associations, because carbonate
79 minerals do not form at acidic pH<6 conditions (Fairén et al., 2004; Fairén, 2013; Peretyazhko et
80 al., 2016; Zolotov and Mironenko, 2016). The carbonate/smectite mineralogical observation
81 might indicate that early Mars was not exclusively neutral-to-alkaline in pH but experienced
3
82 local and perhaps widespread acidic events. As a result, smectite formation on Mars could occur
83 not only under commonly expected neutral/alkaline conditions but also in acidic environments.
84 In terrestrial environments smectites formed under acidic conditions have been reported
85 in acidic saline lakes, seafloor hydrothermal vents and fumarolic areas (Haymon and Kastner,
86 1986; Story et al., 2010; Hynek et al., 2013). Limited laboratory observations have revealed
87 smectite formation through hydrothermal basalt alteration in acidic pH 3-6 environments (Berger
88 et al., 1987; Ghiara et al., 1993; Abdelouas et al., 1997; Dehouck et al., 2014). We have
89 demonstated formation of saponite through alteration of Mars-analogue glass-rich basalt
90 simulant in hydrothermal oxic and anoxic systems buffered by acetic acid (200 °C, pHRT ~4 pH
91 measured at room temperature, (Peretyazhko et al., 2016)).
92 The source of acidity on early Mars remains a subject for discussion. Two major acidity
93 sources have been proposed for early Mars: (1) Fe(II) oxidative hydrolysis (Tosca et al., 2008;
94 Hurowitz et al., 2010), and (2) volcanic release of sulfur dioxide, SO2 (Zolotov and Mironenko,
95 2007; Berger et al., 2009; Gaillard and Scaillet, 2009; Righter et al., 2009; Gaillard et al., 2013;
96 Zolotov and Mironenko, 2016). We have previously hypothesized that sulfuric acid (H2SO4)
97 produced by volcanic SO2 degassing on early Mars was the major source of acidity for the
98 alteration of basaltic materials and subsequent formation of smectite (Peretyazhko et al., 2016).
99 Our hypothesis was supported by modeling and mineralogical observations. For instance, recent
100 detection of smectite minerals co-existing with sulfates might indicate that basalt weathering and
101 smectite formation occurred under acidic conditions caused by sulfuric acid (Farrand et al., 2009;
102 Wray et al., 2011; Cavanagh et al., 2015; Flahaut et al., 2015; Rampe et al., 2017).
103 Thermodynamic modeling of basalt interaction with H2SO4-rich solutions revealed formation of
104 Al-rich phyllosilicates (kaolinite, montmorillonite) under mildly acidic conditions followed by
105 Mg/Fe smectites at higher pH (Zolotov and Mironenko, 2007; Zolotov and Mironenko, 2016)
4
106 and modelling of Mars-like aqueous systems predicted coexistence of smectite and sulfate
107 minerals under mildly acidic pH 4-6 conditions (Fairén, 2013).
108 Formation of smectite through basalt alteration under Mars-relevant acid sulfate
109 conditions has not been experimentally studied and the effect of pH and the nature of forming
110 phyllosilicate minerals remains unknown. The objective of this work was to investigate
111 formation of smectite through hydrothermal alteration of Mars-analogue basalt in the presence of
112 sulfuric acid of variable concentrations, to assess acidity sources and to determine the extent to
113 which H2SO4 is the source of acidity on early Mars.
114
115 2. MATERIALS AND METHODS
116
117 2.1. Synthesis and characterization of Adirondack basalt simulant
118 Synthetic basalt simulant of composition similar to that for Adirondack class rocks
119 analyzed by the Mars Exploration Rover Spirit at Gusev Crater (McSween et al., 2006) was used
120 in smectite formation studies (hereafter, denoted as Adirondack basalt simulant). The detailed
121 synthesis procedure of the Adirondack basalt simulant is reported by Peretyazhko et al., (2016).
122 Briefly, a powdered mixture of reagent-grade oxides and carbonates was melted at 1400 °C in an
123 Au-Pt alloy crucible for 3d under the oxygen fugacity IW+1 (IW = iron-wüstite buffer) and
124 quenched in water. The glassy product was then crushed, ground and sieved to <53 µm particle
125 diameter for characterization and smectite formation experiments. X-ray diffraction analysis
126 revealed that Adirondack simulant contained 66 wt% X-ray amorphous glass, 32 wt% olivine
127 and 2 wt% chromite quench crystals within the glass matrix (Peretyazhko et al., 2016).
128 Composition of the glass phase, olivine and chromite are summarized in Table EA-1.
129
5
130 2.2. Hydrothermal smectite formation experiments and characterization
131 Adirondack basalt simulant suspensions of 17 g/l (water to rock ratio = 60) were prepared
132 by mixing 250 mg simulant with 15 ml of solutions having variable H2SO4 concentrations (11.0
133 ± 0.1 mM, 13.6 ± 0.1 mM, 16.3 ± 0.1 mM, 21.9 ± 0.1 mM, 30.6 ± 0.3 mM and 42.5 ± 0.2 mM).
134 Initial sulfuric acid concentrations were measured as total sulfate by ion chromatography as
135 described below. Duplicate samples were placed in batch reactors (Teflon lined 23 ml Parr acid
136 digestion vessel) and incubated in an oven at 200 °C for 14d. The temperature of 200 °C was
137 chosen because hydrothermal conditions are favorable for smectite formation promoting
138 breaking of chemical bonds and rapid basalt alteration (Berger et al., 2014). However, smectite
139 formation through alteration of basalts has been shown to occur at lower temperatures (Seyfried
140 et al., 1978; McMurtry and Yeh, 1981; Kloprogge et al., 1999; Hynek et al., 2013). Sample
141 preparation and hydrothermal incubation were performed under ambient atmosphere (oxidizing
142 conditions). Our previous study (Peretyazhko et al., 2016) demonstrated that smectite formed
143 through alteration of Adirondack basalt simulant under oxidizing and reducing acidic
144 hydrothermal conditions and that oxidizing conditions only affected Fe oxidation state in
145 smectite.
146 Suspension pH was measured immediately after mixing Adirondack basalt simulant with
147 H2SO4 (pH0d) and after opening the reactors at the end of 14d incubation (pH14d) with AccupHast
148 pH electrode using Orion Star A329 meter. Both measurements were performed at room
149 temperature (RT) of ~25 ºC. After pH measurement, solution was collected in a syringe and
150 passed through a 0.2 µm syringe filter (PVDF, Fisherbrand). An aliquot of each filtered solution
151 was acidified with ultra-pure HNO3 (Fisher) for dissolved Ca, Mg, Na, Si, Al and Fe analysis by
152 Inductively Coupled Plasma Mass Spectrometry (ICP-MS, Element XR Thermo Scientific) with
153 argon as the carrier gas. The remaining solution was analyzed for sulfate by ion chromatography
6
154 using Dionex ICS-2000 equipped with Dionex IonPac AS18 column (4 x 250 mm), Dionex EGC
155 III KOH eluent and a suppressed conductivity detector Dionex AERS 500 4 mm with a 20 µm
156 injection volume.
157 Separated solids were dried in an oven at 40 °C for ~1h prior to analysis. X-ray
158 diffraction (XRD) patterns were recorded using a Panalytical X’Pert Pro with Co Kα radiation.
159 Samples were analyzed at 45 kV/40 mA with a 0.02 2 step for 1 min. Programmable divergent
160 and antiscatter slits were set to 1/4 with manual 1/2 antiscatter slit. The instrument was
161 operated under ambient conditions and calibrated with novaculite (quartz) standard (Gemdugout,
162 State College, PA). Standard powder mounts and low background silicon holders were used for
163 analyses of bulk samples and clay size fractions, respectively. Clay size fractions were separated
164 by ultrasonic dispersion (Kachanoski et al., 1988). Smectite-containing suspensions were
165 ultrasonically treated with a probe-type sonicator (Soniprep 150) for 10s. The suspension was
166 then allowed to sediment by gravity for ~1hr and supernatant containing the clay size fraction
167 was collected. The procedure was repeated until supernatant solutions became clear and the
168 collected clay fraction was separated by centrifugation. The clay fraction slurry was spread over
169 the low background holder with a pipette and dried at RT prior to analysis. Clay fractions were
170 treated with glycerol and KCl in order to confirm formation of smectite (Whittig and Allardice,
171 1986).
172 Visible and near-infrared reflectance (VNIR) spectra (0.35-2.5 µm) were measured with
173 an Analytical Spectral Devices FieldSpec3 fiber-optic based spectrometer with a reflectance
174 probe attachment (internal light source). Spectra were collected in absolute reflectance mode
175 using a Spectralon standard. Continuum removed spectra were obtained using ENVI software
176 (ENvironment for Visualizing Images, Harris Geospatial Solutions).
7
177 Mössbauer analyses were performed at RT under ambient conditions using MIMOS-II
178 instruments that are the laboratory equivalent of the instruments onboard the Mars Exploration
179 Rovers [SPESI, Inc., (Klingelhoefer et al., 2003)]. Velocity calibration was carried out with the
180 program MERView (Agresti et al., 2007) using the MIMOS-II differential signal and the
181 spectrum for metallic Fe foil acquired at RT. Mössbauer parameters [center shift (CS) with
182 respect to metallic Fe foil at RT, quadrupole splitting (QS), magnetic hyperfine field strength
183 (Bhf), and subspectral areas (A)] were obtained with a least-squares fitting procedure using
184 MERFit (Agresti and Gerakines, 2009). Reported values for A were converted to % Fe using the
185 f-factor ratio, fFe(III)/f Fe(II) = 1.21 (Morris et al., 1995). Uncertainty for CS, QS, and full width at
186 half maximum (FWHM) was ± 0.02 mm/s and uncertainty for A was ± 2% absolute. The
187 following constraints were used in the least squares fitting procedure. (1) For each doublet, the
188 component peaks were constrained to have equal widths and areas. (2) Because of the low
189 observed intensity of hematite sextets, values for CS and QS were fixed to literature values
190 (e.g.,(Morris et al., 1985)), peak areas were constrained to 3 : x : 1 : 1 : x : 3, where x is variable
191 and FWHM decreased symmetrically toward zero velocity with skewed Lorentzian lineshapes.
192 Four representative samples were analyzed by Mössbauer spectroscopy covering the whole range
193 of the final pH14d values.
194 Electron microprobe analyses of polished sections of smectite-containing Adirondack
195 basalt simulant were analyzed for elemental composition using a CAMECA SX-100 electron
196 microprobe (Figure EA-1). Backscattered electron imaging and analyses were collected at an
197 accelerating voltage of 15 kV and probe current of 20 nA with a spot size of 1 µm. Element
198 concentrations were calibrated using University of Tokyo Oxides standards. All measurements
199 were made under 20s peak counting and 10s background counting. Because of high water
200 content in smectite, total elemental content was lower than 100% in all analyzed samples. Data
8
201 with totals ≥74 wt% were used for smectite formula calculation. Residual sulfur and
202 phosphorous were considered as impurities and subtracted from the totals in formula
203 calculations. All Si atoms were assigned to tetrahedral sites and the remaining tetrahedral sites
204 were filled with Al (Al = 8-Si). The remaining Al atoms, Mg, Fe and minor Ti, Ni, Cr and Mn
205 were assigned to the octahedral sites, while Na, Ca and K were assigned to interlayer sites (Bain
206 and Smith, 1987).
207
208 2.3. Calculations of sulfur degassing on Mars
209 Crust production model and the amount of sulfur degassing from magma were used to
210 determine the release of SO2 during Noachian/Early Hesperian epoch on Mars and then to obtain
+ 211 the total amount of acidity (total amount of protons, H ) produced from the oxidation of SO2 to
212 H2SO4.
213 First, a crust production model described by Hirschmann and Withers (2008) was used to
214 calculate crust accumulation from early Noachian (~4.5 Ga) to late Amazonian (0 Ga). The
215 model is based on assumptions that crust formation is proportional to the heat flow and that total
216 crust production during 4.5 Ga is 654·106 km3 as estimated by (Greeley and Schneid, 1991). The
217 model is described by the equation:
218 (1)
219 where variables are V (crust production) and t (time) and constants are a = 9.277·105 km3, b =
220 2.385·105 km3, c = 6·10-3 Ma-1, d = 3.98·10-4 Ma-1 (Hirschmann and Withers, 2008). The
221 calculated total crust production is shown in Figure EA-2. The amount of sulfur released from
222 magma was then calculated as
223 Total S= Smagma ρ V (2)
9
3 224 where is an average basalt density of 2800 kg/m and Smagma is the amount of sulfur degassed
225 from magma of 2400 mg/kg (Righter et al., 2009).
226
227 3. RESULTS
228
229 3.1. pH before and after incubation
230 Different initial H2SO4 concentrations led to variable degrees of basalt neutralization
231 after 14d incubation. The average initial pH for all experiments was acidic (pH0d ≤ 2), and the
232 final pH14d varied over a wide range including acidic (pH14d 1.8 and 2.2), moderately acidic
233 (pH14d 3.1 and 3.7), neutral (pH14d 7.2) and slightly alkaline (pH14d 8.4) (Figure 1).
234 3.2. Mineralogical characterization
235 3.2.1. X-ray diffraction
236 Powder XRD analyses of reaction products showed increasing phyllosilicate formation at
237 pH14d ≥3 on the basis of increasing intensity of the 001 diffraction peak at ~15 Å while no
238 phyllosilicate formation occurred at pH14d 1.8 and 2.2 (Figure 2a, Table 1). After separation of
239 the clay size fraction, the phyllosilicate was identified as smectite based on expansion of the 001
240 peak to 17.4-18 Å after glycerol treatment, its collapse to 12.8 Å at room temperature upon KCl
241 addition, and its additional collapse to 10.1 Å after heating at 550 °C (Figure EA-3). The
242 positions of 02l and 060 d-spacings allowed distinguishing dioctahedral from trioctahedral
243 smectite (Moore and Reynolds, 1997; Vaniman et al., 2014). The 02l and 060 d-spacings
244 revealed that trioctahedral smectite formed at pH14d 3.7, 7.2 and 8.4 and the peak positions
245 (Figure 2b, Table 2) were within the range reported for saponite (02l: 4.52-4.64 Å and 060: 1.53-
246 1.56 Å (Treiman et al., 2014; Chemtob et al., 2015; Peretyazhko et al., 2016)). The 02l and 060
247 d-spacings at 4.48 Å and 1.50 Å, respectively, indicated formation of dioctahedral smectite at
10
248 pH14d 3.1 (Figure 2b, Table 2). The peak positions were in agreement with d-spacing reported for
249 montmorillonite (02l: 4.47 Å and 060: 1.492-1.504 Å, (Moore and Reynolds, 1997; Vaniman et
250 al., 2014)). However, formation of montmorillonite could not be confirmed with VNIR and
251 electron microprobe analyses (sections 3.2.3 and 3.2.4) and therefore phyllosilicate formed in the
252 pH14d 3.1 sample is referred as dioctahedral smectite. Formation of smectite in all samples was
253 accompanied by decrease in intensity of the broad basaltic glass peak between 20° and 40° 2θ
254 with respect to the unaltered material (Figure 2a) suggesting glass as the smectite progenitor.
255 Anhydrite (CaSO4) was found to precipitate at all values of pH14d and natroalunite
256 (NaAl3(SO4)2(OH)6) was detected only at pH14d 1.8 (Figure 2a, Table 1). A weak diffraction
257 peak of hematite (α-Fe2O3) was observed in the pH14d 1.8 sample (Figure 2a, Table 1), and the
258 presence of hematite was confirmed by Mössbauer spectroscopy (Section 3.2.2). Hematite was
259 not detected in any other samples by XRD, but Mössbauer analysis showed that hematite was
260 present in the pH14d range from 1.8 to 8.4 (Section 3.2.2). Chromite, present in the unaltered
261 material, did not dissolve during incubation while olivine completely dissolved in the pH14d 1.8
262 and 2.2 samples as evident from the disappearance of its sharp diffraction peaks (Figure 2a,
263 Table 1).
264 3.2.2. Mössbauer spectroscopy
265 Mössbauer analysis of unaltered Adirondack basalt simulant showed that Fe was
266 associated with olivine and glass phases, 19 ± 2 and 81 ± 2% respectively and was mostly Fe(II)
267 (Table EA-2, (Peretyazhko et al., 2016)).
268 Analysis of altered Adirondack basalt simulant revealed formation of smectite through
269 alteration of the glass phase. The decrease in the concentration of total Fe associated with the
270 glass phase (39 ± 2% pH14d 3.1, 43 ± 2% pH14d 3.7 and 54 ± 2% pH14d 8.4, Figure 3a-c, Table
271 EA-2) compared to the unaltered simulant, was equal within error to amount of total Fe
11
272 associated with the smectite formed after 14d incubation (38 ± 2% pH14d 3.1, 41 ± 2% pH14d 3.7
273 and 56 ± 2% pH14d 8.4, Figure 3a-c, Table EA-2). Smectite spectra were fit with one Fe(II)
274 doublet (CS = 1.04 -1.09 mm/s and QS = 2.42 -2.66 mm/s) and one Fe(III) doublet (CS = 0.41 -
275 0.48 mm/s and QS = 0.54 - 0.70 mm/s QS, Table EA-2) and assigned to Fe(II) and Fe(III) in
276 octahedral coordination (Treiman et al., 2014). The smectite was oxidized with Fe(III)/Fe equal
277 to 0.61, 0.73 and 0.82 at pH14d 3.1, 3.7 and 8.4, respectively.
278 The most acidic pH14d 1.8 sample was also characterized by Fe(II) and Fe(III) doublets
279 with parameters similar to those assigned to smectite in the other samples (Fe(II) doublet: CS =
280 1.03 mm/s and QS = 2.60 mm/s and Fe(III) doublet: CS = 0.46 mm/s and QS = 0.62 mm/s,
281 Table EA-2). However, smectite did not form in this sample within XRD detection limits (Table
282 1). The Fe might be associated with an altered glass layer which forms during first stage of acidic
283 basalt alteration (Berger et al., 1987; Oelkers, 2001) and/or with an incipient phyllosilicate with
284 domains too small to coherently scatter X-rays.
285 The concentration of total Fe associated with olivine progressively decreased with
286 decreasing pH14d, and no Fe in olivine was detected in the pH14d 1.8 sample as shown in Table
287 EA-2. Olivine dissolution by acid-sulfate solutions could contribute Fe to smectite precipitation.
288 However, the Mössbauer evidence is that Fe(II) from olivine dissolution was oxidized and
289 precipitated as hematite, because the decrease in concentration of Fe(II) associated with olivine
290 was associated with a concomitant increase in the concentration of Fe(III) associated with
291 hematite (Figure 3, Table EA-2). Iron content in hematite (15 ± 2% pH14d 1.8, 12 ± 2% pH14d
292 3.1, 8 ± 2% pH14d 3.7 and 2 ± 2% pH14d 8.4, Figure 3, Table EA-2), was equal within an error of
293 fitting to decrease in olivine Fe content with respect to the unaltered Adirondack basalt simulant
294 (19 ± 2% pH14d 1.8, 11 ± 3% pH14d 3.1, 6 ± 3% pH14d 3.7 and 4 ± 3% pH14d 8.4, Figure 3, Table
12
295 EA-2). The hematite peaks were broad and skewed toward zero velocity, implying poor
296 crystallinity (small particle diameter) and potentially Al substitution for Fe(III).
297 3.2.3. Smectite composition and structural formula
298 The structural formula calculated from electron microprobe analysis of smectite-containing
299 particles (Table EA-3, Figure EA-1) showed that smectite contained Si and Al in tetrahedral
300 layers, Al, Mg, Fe with traces of Ti, Mn, Ni and Cr in octahedral layers, and Na, Ca and K
301 occupied interlayer sites. Given the high uncertainties in the measured data due to high water
302 content in smectite, the results of electron microprobe analysis did not allow to accurately
303 quantify smectite composition. Electron microprobe analysis indicated that smectite enriched in
304 Fe and Mg formed at pH14d 3.7, 7.2 and 8.4 while smectite equally enriched in Al, Mg and Fe
305 formed at pH14d 3.1 (Table EA-3). The average total number of octahedral cations in a unit cell
306 was close to 6 in the pH14d 3.7, 7.2 and 8.4 samples (Table EA-3) and was consistent with
307 trioctahedral saponite. The total number of octahedral cations was close to 5 in the pH14d 3.1
308 sample (Table EA-3) which might indicate formation of smectite containing dioctahedral and
309 trioctahedral domains.
310 3.2.4. Visible and near-infrared spectroscopy
311 VNIR spectral features between ~1.2 and 2.5 µm (Figure 4, Table 2) for saponite formed
312 at pH14d 3.7, 7.2, and 8.4 were assigned to OH in H2O and M-OH (~1.4 µm), to interlayer and
313 adsorbed H2O (~1.9 µm), and to trioctahedral M3-OH and dioctahedral M2-OH functional groups
314 (2.2-2.5 µm), where M is any combination of Al, Mg, Fe(II) and Fe(III) as octahedral cations.
315 The band centered between 2.30 and 2.32 μm was assigned to (Mg, Fe(II), Fe(III))3OH
316 combination stretching and bending vibration (Neumann et al., 2011; Treiman et al., 2014). (Mg,
317 Fe(II), Fe(III))3OH band center position was shifted to shorter wavelength in the pH14d 8.4
318 sample with respect to the pH14d 3.7 and 7.2 samples (Figure 4b, Table 2). The band position
13
319 shift could be due to increase in degree of Fe oxidation in saponite from Fe(III)/Fe = 0.73 at
320 pH14d 3.7 to Fe(III)/Fe = 0.82 at pH14d 8.4, (Table EA-2, (Peretyazhko et al., 2016)). The
321 position variability for the ~2.31 μm band might also result from heterogeneous distributions of
322 cations on octahedral sites [e.g., (Fe(III)2Mg)-OH versus (Fe(III)Mg2)-OH]. Saponite observed at
323 pH14d 7.2 and 8.4 had a weaker band near 2.40 μm band commonly observed in Fe/Mg smectites
324 (e.g., Carter et al., 2015; Chemtob et al., 2015). The exact band assignment is uncertain but
325 might be due to (Mg, Fe(II), Fe(III))3OH combination stretching and translation vibration (Frost
326 et al., 2002; Chemtob et al., 2015). The spectral feature near 2.24 μm was assigned to Al(Fe(III),
327 Fe(II), Mg)OH (Madejová et al., 2011; Chemtob et al., 2015) and indicated the presence of
328 octahedral Al, possibly present as dioctahedral domains as previously observed in synthetic
329 saponite (Chemtob et al., 2015; Peretyazhko et al., 2016). Spectral properties of saponite
330 synthesized at pH14d 3.7, 7.2 and 8.4 were similar to the properties reported for smectite on Mars
331 by VNIR spectrometers. The bands near 2.30 and 2.40 µm has been reported in Fe/Mg-smectite
332 identified on Mars while the band at 2.24 µm is present in about 10% of Fe/Mg-phyllosilicate
333 spectra (Carter et al., 2013), including nontronite (Bishop et al., 2013) but not saponite.
334 The spectrum of the pH14d 3.1 sample containing dioctahedral smectite was similar to the
335 spectra of the untreated Adirondack basalt simulant and pH14d 1.8 and 2.2 samples (Figure 4b).
336 All these samples had a broad band centered on 2.21-2.22 µm likely resulting from vibration of
337 Si-OH associated with the unaltered glass phase. The diagnostic band of montmorillonite at 2.20
338 µm was not distinguishable from the basalt band because of relatively low smectite abundance at
339 pH14d 3.1 (indicative by low intensity 001 diffraction peak in the sample, Figure 2a).
340 Hematite presence was evident from the ferric absorption edge between ~0.50 um and
341 ~0.75 um (Figure 4a) corresponding to the onset of strong absorption and a relative reflectivity
342 maximum for hematite (Morris et al., 1985). The VNIR manifestation of the hematite was most
14
343 noticeable for the pH14d 1.8 sample which had the highest hematite content (15% of total Fe,
344 Table EA-2)
345 3.3. Characterization of aqueous phase
346 Mineralogical transformations of Adirondack basalt simulant to smectite, anhydrite,
347 hematite and natroalunite (Table 1) resulted in release of variable proportions of Si, Mg, Na and
348 Ca. Complete dissolution of olivine in the pH14d 1.8 and 2.2 samples and subsequent release of
349 Mg and Si into solution (Figure 5a) were lower than the expected Si and Mg concentrations if all
350 olivine was dissolved (32 ± 3 mM Si and 42 ± 4 mM Mg based on the total SiO2 and MgO
351 contents in olivine, Table EA-1). The results suggest that Si and Mg dissolved from olivine were
352 partially incorporated into an altered glass phase under acidic conditions (Berger et al., 1987).
353 Smectite formation in the pH14d range from 3.1 to 8.4 (Figure 2a) led to decrease in Si
354 and Mg dissolution (Figure 5a) where the Mg dissolution likely controlled the nature of the
355 forming smectite mineral (Meunier, 2005). Dioctahedral smectite was only detected in the pH14d
356 3.1 sample in which nearly 50% of the total Mg was leached into solution (calculations are based
357 on dissolved Mg concentrations in Figure 5a and total MgO in the glass phase, Table EA-1)
358 while formation of trioctahedral saponite (pH14d 3.7, 7.2 and 8.4) was accompanied by lower or
359 no Mg release into solution.
360 Dissolved Na was affected by precipitation of natroalunite and smectite formation.
361 Dissolved Na concentration was low and did not change in the pH14d range from 1.8 to 3.1 then
362 increased as pH shifted from acidic to slightly alkaline (Figure 5b). Low release of Na from
363 basaltic glass at pH14d 1.8-3.1 could result from precipitation of natroalunite, although the
364 presence of this phase was only confirmed at pH14d 1.8 (Table 1). Dissolved Na increased in the
365 pH14d range from 3.1 to 8.4 (Figure 5b) due to Na loss from glass phase during alteration of
366 Adirondack basalt simulant into smectite (Peretyazhko et al., 2016). Similar to Na, Ca was
15
367 expected to increase in solution with increasing basalt alteration as pH rose (Berger et al., 1987).
368 However, dissolved Ca concentration did not change (Figure 5b), and such non-pH-dependent
369 Ca release, together with the observed decrease in sulfate concentration (Figure 5c), resulted
370 from anhydrite precipitation (Figure 2a, Table 1) which is favorable at elevated temperatures
371 (Bishop et al., 2014).
372 Solution concentrations of Al and Fe were low for all samples resulting from their
373 incorporation into secondary phases (Figure EA-4, Tables 1 and 2). Dissolved Al did not exceed
374 0.04 mM (Figure EA-4a) as a result of precipitation of natroalunite at pH14d 1.8, low release from
375 glass alteration and incorporation into smectite at higher pH values (Tables 2 and EA-3) and
376 potentially in hematite. Dissolved Fe was <0.15 mM (Figure EA-4b) in all samples due to
377 precipitation of hematite accompanying olivine dissolution (Figure 3), low release from altering
378 glass phase and smectite formation (Tables 2 and EA-3).
379
380 4. DISCUSSION
381
382 The formation of smectite in the Noachian and early Hesperian has been attributed to
383 aqueous alteration of basaltic materials under neutral to alkaline pH aqueous conditions (Bibring
384 et al., 2006; Chevrier et al., 2007; Ehlmann et al., 2009; Bishop et al., 2013; Bristow et al., 2015;
385 Flahaut et al., 2015; Jain and Chauhan, 2015). However, the results of our experiments show that
386 formation of smectite from Mars-analogue basalt simulant is possible in much more acidic
387 environments. We found that gradual acid neutralization during acid-sulfate alteration of the
388 basalt simulant under hydrothermal conditions led to smectite formation at pH14d ≥ 3. The nature
389 of the forming smectite mineral varied with the solution final pH. Dioctahedral smectite equally
16
390 enriched in Al, Fe and Mg formed at pH14d ~3 followed by saponite enriched in Fe and Mg at
391 pH14d close to 4 or higher.
392 Our experimental results demonstrate that smectite formation through basalt alteration on
393 Mars was not limited to neutral/alkaline pH conditions but could partially occur under acidic
394 conditions once solution pH was close to 3 or higher. Sources of acidity on early Mars and the
395 persistence of acidic aqueous environments have been controversial. The basaltic crust should
396 provide enormous acid neutralizing potential (Niles and Michalski, 2009) suggesting that
397 aqueous systems on Mars should be largely alkaline. However, as noted above, absence of large
398 scale carbonate deposits shows an evidence for acidic conditions even in heavily altered regions
399 of the martian crust (Fairén et al., 2004). The principal acidity sources proposed for early Mars
400 are Fe(II) oxidative hydrolysis (Tosca et al., 2008; Hurowitz et al., 2010) and SO2 degassing
401 (Berger et al., 2009; Gaillard and Scaillet, 2009; Righter et al., 2009; Gaillard et al., 2013). We
402 argue here that SO2 degassing was the major source of acidity on early Mars generating H2SO4
403 sufficient for acidic phyllosilicate formation through basalt alteration and Fe(II) oxidative
404 hydrolysis (i.e., Fe(III) hydrolysis) was likely not the source of wide-spread acidity on Mars.
405 4.1. Acidity sources on early Mars: Fe(II) oxidative hydrolysis
406 Iron(III) hydrolysis was first suggested as the source of acidity at Meridiani Planum
407 (Hurowitz et al., 2010) and was further generalized as a possible widespread acidity source on
408 early Mars (e.g., (Ehlmann et al., 2011; Bowen et al., 2012; Horgan and Bell III, 2012;
409 McLennan, 2012; Marlow et al., 2014; Kaplan et al., 2016)). The Fe(II) oxidative hydrolysis
410 model proposed that upwelling of anoxic Fe(II)-containing groundwater generated from basalt
411 dissolution under anoxic circum-neutral pH conditions could lead to surface oxidation of Fe(II)
412 to Fe(III) followed by Fe(III) hydrolysis, release of H+ and precipitation of Fe(III) secondary
17
413 minerals (Hurowitz et al., 2010). The latter include nanophase iron(III) (hydr)oxides (interpreted
414 as schwertmannite (FeO(OH)0.75(SO4)0.125) (Hurowitz et al., 2010)), jarosite
415 ((K,Na)Fe3(SO4)2(OH)6) and hematite (α-Fe2O3) detected by Mössbauer spectroscopy at
416 Meridiani Planum (Morris et al., 2006). The reactions below summarize the mechanism
417 proposed for acidity generation following the sequence developed by Hurowitz et al., (2010)
418 with the exception that eq. 3 is included for basalt dissolution (eq. 3 written for olivine Fe(II)-end
419 member fayalite (Fe2SiO4)):
420 + 2+ 421 0.5Fe2SiO4 + 2H → Fe + 0.5SiO2 + H2O or (3a)
2+ - 422 0.5Fe2SiO4 + H2O → Fe + 0.5SiO2 + 2OH (3b)
2+ + 3+ 423 Fe + 0.25O2 + H → Fe +0.5H2O (4)
3+ 2- + 424 Fe + 0.125SO4 + 1.75H2O → FeO(OH)0.75(SO4)0.125 + 2.75H (5)
+ 2- + + 425 FeO(OH)0.75(SO4)0.125 + 0.25H2O + 0.75H + 0.5417SO4 + 0.33(K or Na ) →
426 0.33(K,Na)Fe3(SO4)2(OH)6 (6)
+ 2- 427 FeO(OH)0.75(SO4)0.125 + 0.25H2O → FeOOH + 0.25H + 0.125SO4 (7)
428 FeOOH → 0.5Fe2O3 + 0.5H2O (8)
429 While groundwater interaction with basalt was suggested to be a source of dissolved
430 Fe(II), basalt dissolution (eq. 3) was not included into Fe(II) oxidative hydrolysis model
431 (Hurowitz et al., 2010). As a result, the model predicted that sufficient amount of H+ was
432 produced through Fe(III) hydrolysis to develop acidic conditions (Hurowitz et al., 2010). If
433 basalt dissolution reactions are added to the model, then no acidity would be generated because
434 Fe(II) release is accompanied by OH- release (e.g., eq. 3b, Text EA-1). While our calculations
435 only account for Fe, the release of Mg, Ca, and Na from the basalts will also be accompanied by
436 release of OH- which will contribute to generation of additional alkalinity.
18
437 The OH- generated from dissolution of basalt (eq. 3) could be neutralized by H+ produced
438 through oxidative dissolution of Fe(II) sulfides present in basalt. However, the produced H+ will
439 not be sufficient to neutralize all released OH- on the basis of current estimates of low
440 abundances of Fe(II) sulfides in martian basalts (1600 mg/kg S in shergottite sulfides or 0.16
441 wt%, (Righter et al., 2009)). Alternatively, the OH- released from basalt together with Fe(II),
442 Mg, Ca and Na could be neutralized if the anoxic alkaline solutions encounter anoxic acidic
443 solutions. If all released OH- is neutralized, then Fe(II) oxidation followed by hydrolysis under
444 oxic conditions will lead to development of acidic conditions according to the Hurowitz et al.,
445 (2010) model (eqs. 4-8). However, such scenario will require presence of acidic conditions, for
446 instance H2SO4, before Fe(III) hydrolysis occurred.
447
448 4.2. Acidity sources on early Mars: SO2 degassing and H2SO4 formation
449 Magmatic sulfur degassing was the dominant source for sulfates deposited on the surface
450 of Mars (Righter et al., 2009), but was SO2 release on early Mars able to generate the amount of
451 H2SO4 sufficient for acidic phyllosilicate formation through alteration of basalt?
452 The calculated amount of sulfur release (eq. 2) during Noachian/early Hesperian epoch
453 was 2.5 x1021 g [~4.5 Ga - ~3.4 Ga; an intermediate value of 3.4 Ga was used for early
454 Hesperian/Late Hesperian boundary which is around 3.2 Ga - 3.6 Ga (Hartmann, 2005); Figure
20 + 455 6]. If all sulfur was released as SO2 and converted to H2SO4 then around 1.6x10 moles of H
456 was produced during dissociation of all acid in water. Formation of sulfuric acid might occur
457 through disproportionation of SO2 to H2SO4 and elemental S and S could be further transformed
458 into H2SO4 in the presence of oxidizing agent and water (Habashi and Bauer, 1966; Halevy et al.,
459 2007). It should be noted, however, an assumption that all degassed sulfur was transformed into
19
460 sulfuric acid likely overestimates the amount of produced acidity but our calculations allows to
461 potentially evaluate an upper limit of acidity that could have been generated on early Mars.
462 The protons generated from sulfuric acid are neutralized by Ca, Mg and Na in basalt via
463 proton exchange leading first to formation of an altered hydrated silica layer enriched in Fe and
464 Al (Berger et al., 1987; Peretyazhko et al., 2016) by schematic reactions:
465 Na-basalt + H+ = H-basalt + Na+ (9)
466 Ca-basalt + 2H+ = 2H-basalt + Ca2+ (10)
467 Mg-basalt + 2H+ = 2H-basalt + Mg2+ (11)
468 Crystallization of phyllosilicate from an altered layer does not substantially change pH as shown
469 in the previous studies of smectite formation through basalt alteration (Ghiara et al., 1993; De La
470 Fuente et al., 2002; Peretyazhko et al., 2016). Iron was not included into calculations because
471 overall process including dissolution of olivine followed by Fe(II) oxidation, Fe(III) hydrolysis
472 and hematite precipitation (Table 1) does not produce or consume H+ (Text EA-1).
+ 473 Approximately 9.4 molH+/kgbasalt H is required for complete neutralization of Na, Ca and Mg in
474 basalt of Adirondack composition (Text EA-2). The amount of H+ generated from volcanic
20 475 outgassing of SO2 (1.6x10 moles) during the Noachian and Hesperian would be sufficient for
476 neutralization of Ca, Mg and Na along with phyllosilicate formation in a 42 m thick basalt layer
477 over the entire planet (Text EA-3). Given that formation of smectite requires Mg to be present in
478 the altering basalt (Kloprogge et al., 1999), partial H+ neutralization by cations would be a more
479 realistic scenario on Mars. Noachian-aged terrains such as Mawrth Vallis have been estimated to
480 contain 20-65% smectite (Poulet et al., 2008). Assuming that the smectite abundance in Mawrth
481 Vallis is representative of H+ neutralization and, consequently, degree of basalt alteration in
482 ancient martian terrains, the calculated average global depth of basaltic alteration would be
483 between 60 m (65% smectite content) to 200 m (20% smectite content) on the entire planet under
20
484 the constraint that alteration of 42 m thick basalt leads to 100% neutralization of Na, Ca and Mg.
485 The calculated depth range is consistent with phyllosilicate layer thickness of a few hundreds of
486 meters estimated in Noachian-aged terrains (Loizeau et al., 2012). Our acidity calculations and
487 experimental results, therefore, provide evidence that SO2 degassing would result in production
488 of large amounts of H+ and gradual neutralization of sulfuric acid during basalt alteration could
489 be an important mechanism of smectite formation on early Mars.
490
491 5. SIGNIFICANCE AND CONCLUSIONS
492
493 Our experiments demonstrated formation of smectite through hydrothermal acid-sulfate
494 alteration of Mars-analogue glass-rich basalt (66 wt% glass, 32 wt% olivine and 2 wt%
495 chromite). Basalt simulant suspensions were acidic with initial pH0d ≤ 2 while final pH14d varied
496 from acidic to mildly alkaline (pH14d 1.8-8.4) at the end of 14d incubation at 200 ºC. Alteration
497 of the glass phase of basalt simulant resulted in formation of dioctahedral smectite as pH14d
498 reached ~3 followed by trioctahedral saponite at pH14d close to 4 or higher. Reaction of
499 Adirondack basalt simulant with sulfuric acid also led to formation of anhydrite at all pH14d
500 values while natroalunite was detected at pH14d 1.8. Hematite precipitated as a result of partial or
501 complete oxidative dissolution of olivine in the pH14d range from 1.8 to 8.4.
502 Acidity on early Mars can be accounted for by volcanic outgassing of SO2 and formation
503 of sulfuric acid. Acid-sulfate conditions, suitable enough to form smectite and prevent carbonate
504 formation were likely sustained by low water/rock ratio and short duration of water-rock
505 interaction (Berger et al., 2009; Berger et al., 2014). Acidic smectite formation areas on early
506 Mars might be constrained to near-surface hydrothermal areas with magmatic outgassing
507 resulted from volcanic or impact activity (Solomon et al., 2005; Hynek et al., 2013). In these
21
508 systems only small portions of crust are exposed to acidity generated by H2SO4 formed from
509 degassed SO2. Short-term aquatic systems analogous to ephemeral acid saline lakes in Western
510 Australia might also allow development of acidic conditions (Story et al., 2010; Bowen et al.,
511 2012). Although there may also be acidity generated by the Fe(II) oxidative hydrolysis in local
512 environments, our calculations indicate that this mechanism is not a plausible large-scale acidity
513 source because of the alkalinity that is generated during basalt dissolution.
514 Formation of phyllosilicate by sulfuric acid alteration of basaltic materials on Mars would
515 likely be accompanied by precipitation of sulfate minerals including anhydrite. Although
516 anhydrite cannot be detected by the same remote sensing measurements used to detect
517 phyllosilicates because of the absence of diagnostic H2O spectral features in the VNIR, anhydrite
518 might coprecipitate with phyllosilicate in deposits formed under hydrothermal sulfuric acid
519 environments (>60 °C (Bishop et al., 2014)). Acid-sulfate smectite formation might also occur in
520 younger terrains. Phyllosilicates (e.g., smectite) co-existing with anhydrite in Gale crater may
521 have formed along with the anhydrite under acid sulfate conditions (Rampe et al., 2017). The
522 hypothesis is supported by occurrence of jarosite in these mudstones in Gale crater indicative of
523 acidic environments (Rampe et al., 2017). Fe/Mg-rich phyllosilicates interbedded with sulfate
524 minerals in Meridiani Planum have been proposed to form prior to sulfates in moderate pH,
525 sulfate-free aqueous environments followed by sulfate precipitation from acidic surface and
526 groundwater (Flahaut et al., 2015). Our findings demonstrate that multiple events are not
527 required to form interbedded phyllosilicates and sulfates deposits. Acid-sulfate alteration can
528 explain martian mineralogical observations including mixed phyllosilicate/sulfate deposits and
529 lack of abundant carbonate deposits.
530
531
22
532 ACKNOWLEDGEMENTS
533
534 We thank Z. Peng for performing ICP-MS analysis. We thank Dr. Cuardos, Dr. Zolotov and two
535 anonymous reviewers for valuable suggestions and comments that help to improve the quality of
536 the manuscript. We thank the Associate Editor Dr. Catalano for handling the manuscript.This
537 work was supported by NASA’s Mars Fundamental Research Program grant # 11-MFRP11-
538 0090.
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23
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766
767
768
35
769 Table 1. Phases determined in the 14d-incubated Adirondack basalt simulants by XRD and
770 Mössbauer spectroscopy. Olivine and chromite are unaltered phases from the original basalt
771 simulant.
pH14d Crystalline phases 1.81 ± 0.01 anhydrite, natroalunite, hematitea, chromite 2.19 ± 0.01 anhydrite, chromite 3.12 ± 0.03 dioctahedral smectite, anhydrite, hematite, olivine, chromite 3.72 ± 0.34 saponite, anhydrite, hematite, olivine, chromite 7.17 ± 0.38 saponite, anhydrite, olivine, chromite 8.38 ± 0.20 saponite, anhydrite, hematite, olivine, chromite a 772 Hematite was detected by both XRD and Mössbauer in the pH14d 1.8 sample and by Mössbauer 773 spectroscopy alone in the other samples. Hematite was present within the pH14d range from 1.8 to 774 8.4 and likely precipitated in the pH14d 2.2 and 7.2 samples not analyzed by Mössbauer 775 spectroscopy. 776
777
778
779
780
781
782
783
784
785
786
787
788
36
789
790 Table 2. Smectite characterization data obtained by XRD and VNIR. pH14d Smectite XRD VNIR d-spacing, Å combination band position, µm 02l 060 Al(Fe(III), (Mg, Fe(II), (Mg, Fe(II), b b c Fe(II), Mg)OH Fe(III))3OH Fe(III))3OH dioctahedral 3.12 ± 0.03 4.48 1.50 nda nd nd smectite 3.72 ± 0.34 saponite 4.55 1.53 2.239 2.319 nd 7.17 ± 0.38 saponite 4.58 1.53 2.246 2.319 2.403 8.38 ± 0.20 saponite 4.57 1.53 2.243 2.300 2.408 791 a nd-not detected; b Combination stretching and bending vibration (Neumann et al., 2011; 792 Treiman et al., 2014); c Combination stretching and translation vibration (Frost et al., 2002; 793 Chemtob et al., 2015). 794
795
796
797
798
37
pH0d 8 pH14d
6
pH
4
2
10 20 30 40 Initial sulfate concentration, mM 799 800 801 Figure 1. Initial (pH0d) and final (pH14d) pH in Adirondack basalt simulant suspensions as a 802 function of initial sulfate concentrations. 803
38
804 805 Figure 2. (a) XRD for unaltered and 14d-incubated Adirondack basalt simulant samples with 806 different pH14d. S = smectite (001, 02l and 060 diffraction peaks are marked next to or above the 807 peaks), A = anhydrite, Ol = olivine, C = chromite, NA = natroalunite, H = hematite; (b) 02l and 808 060 diffraction bands for clay fractions separated from the pH14d 8.4 and 3.1 samples. The bands 809 were analyzed with a 0.02 2 step for 1 min and 0.02 2 step for 17 min in the pH14d 8.4 and 810 3.1 clay fractions, respectively.
39
811
812 Figure 3. Mössbauer spectra and fit subspectra for 14d-incubated (a) pH14d 8.4, (b) pH14d 3.7, (c) pH14d 3.1 and (d) pH14d 1.8 Adirondack 813 basalt simulant collected at RT [subspectral areas in % are for Ol = olivine; Total Gl = total glass, a sum of 2D1 and 2D2 doublets 814 (Table EA-2), Total Sme = total smectite, and Total Alt Gl = total altered glass, a sum of 3D2 and 2D4 doublets (Table EA-2)]. TC = 815 total counts, BC = baseline counts.
816 817 Figure 4. (a) VNIR reflectance spectra of unaltered and 14d-incubated Adirondack basalt 818 simulant samples with different pH14d. Bands at 1.4 and 1.9 µm in the saponite-containing pH14d 819 8.4, 7.2 and 3.7 samples are marked with dotted lines. (b) 2-2.5 µm range of continuum removed 820 reflectance spectra. Positions of M3-OH and M2-OH combination bands in the saponite- 821 containing samples are marked with dotted lines (M is any combination of Al, Mg, Fe(II) and 822 Fe(III) as octahedral cations). Characteristic Al2OH band of montmorillonite at 2.21 µm in the 823 pH14d 3.1 sample was not detected. 824 825
826 Figure 5. Concentrations of dissolved (a) Mg and Si, (b) Na and Ca and (c) sulfate in Adirondack 827 basalt simulant suspensions as a function of pH14d. 828
42
829 830 Figure 6. Sulfur degassed on Mars from 4.5 Ga to present calculated using the crust production 831 model (eq. 2, Hirschmann and Withers 2008) and amount of sulfur degassed from 1 kg magma 832 (2400 mg/kg, Righter et al., 2009). The calculated total degassed sulfur from 4.5 Ga to present 833 was 4.4·1021 g and within the range of estimates of total sulfur release on Mars: 2.2·1022 g 834 (McLennan, 2012), 5.4·1021 g (Gaillard and Scaillet, 2009), 1.7·1020 g (Craddock and Greeley, 835 2009) and 4.5·1019 g (Righter et al., 2009). Dashed line shows a boundary between early and late 836 Hesperian. 837
43